Insights from elastic thermobarometry into exhumation of …...Ep2 crystals range from ~ 50 - 300...

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Insights from elastic thermobarometry into exhumation of high-pressure metamorphic rocks from Syros, Greece Miguel Cisneros1,2*, Jaime D. Barnes1, Whitney M. Behr1,2, Alissa J. Kotowski1,3*, Daniel F. Stockli1, and Konstantinos Soukis4

1Department of Geological Sciences, Jackson School of Geosciences, University of Texas at Austin, Austin, TX, USA 5 2*Current address: Geological Institute, ETH Zürich, Zürich, Switzerland 3*Current address: Department of Earth and Planetary Sciences, McGill, Montreal, Canada 4Faculty of Geology and Geoenvironment, NKUA, Athens, Greece

Correspondence to: Miguel Cisneros ([email protected])

Abstract. We combine elastic thermobarometry with oxygen isotope thermometry to quantify the pressure-temperature (P-T) 10

evolution of retrograde metamorphic rocks of the Cycladic Blueschist Unit (CBU), an exhumed subduction complex exposed

on Syros, Greece. We employ quartz-in-garnet and quartz-in-epidote barometry to constrain pressures of garnet and epidote

growth near peak subduction conditions and during exhumation, respectively. Oxygen isotope thermometry of quartz and

calcite within boudin necks was used to estimate temperatures during exhumation and to refine pressure estimates. Three

distinct pressure groups are related to different metamorphic events and fabrics: high-pressure garnet growth at ~1.4 - 1.7 GPa 15

between 500 - 550 °C, retrograde epidote growth at ~1.3 – 1.5 GPa between 400 - 500 °C, and a second stage of retrograde

epidote growth at ~1.0 GPa and 400 °C. These results are consistent with different stages of deformation inferred from field

and microstructural observations, recording prograde subduction to blueschist-eclogite facies and subsequent retrogression

under blueschist-greenschist facies conditions. Our new results indicate that the CBU experienced cooling during

decompression after reaching maximum high-pressure/low-temperature conditions. These P-T conditions and structural 20

observations are consistent with exhumation and cooling within the subduction channel in proximity to the refrigerating

subducting plate, prior to Miocene core-complex formation. This study also illustrates the potential of using elastic

thermobarometry in combination with structural and microstructural constraints, to better understand the P-T-deformation

conditions of retrograde mineral growth in HP/LT metamorphic terranes.

1 Introduction 25

Constraining the pressure-temperature (P-T) evolution of metamorphic rocks is fundamental for understanding the

mechanics, timescales, and thermal conditions of plate tectonic processes operating on Earth. Historically, one of the most

challenging aspects of thermobarometry has been deciphering the P-T evolution of rocks during their exhumation from peak

depths back to the surface (e.g., Kohn and Spear, 2000; Spear and Pattison, 2017; Spear and Selverstone, 1983). Exhumation

P-T paths are particularly challenging to reconstruct because during retrogression rocks are cooled, fluids are consumed by 30

https://doi.org/10.5194/se-2020-154Preprint. Discussion started: 14 September 2020c© Author(s) 2020. CC BY 4.0 License.

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metamorphic reactions, and strain is progressively localized, all of which result in more sluggish reaction kinetics and lesser

degrees of chemical equilibrium (e.g., Baxter, 2003; Carlson, 2002; Jamtveit et al., 2016; Rubie, 1998). These issues are

especially pronounced in high-pressure/low-temperature (HP/LT) environments characteristic of subduction zones.

Elastic thermobarometry offers an alternative to conventional thermobarometry. Rather than relying on equilibrium

metamorphic reactions, this approach constrains the P-T conditions at which a host crystal entraps an inclusion (e.g., Adams 35

et al., 1975a, 1975b; Rosenfeld, 1969; Rosenfeld and Chase, 1961). Because inclusion-host-pair bulk moduli and thermal

expansivities commonly differ, upon ascent, an inclusion develops residual strain(s) that can be determined from measurements

of Raman shifts. A residual inclusion pressure can be calculated from strain(s) by using Grüneisen tensors (Angel et al., 2019;

Murri et al., 2018, 2019) or experimental hydrostatic calibrations (e.g., Ashley et al., 2014; Enami et al., 2007; Thomas and

Spear, 2018). Elastic modeling is then used to calculate the initial entrapment conditions of when the host grew around the 40

inclusion, and thus can be used to determine the conditions at which individual host minerals grew during metamorphism (e.g.,

Alvaro et al., 2020; Ashley et al., 2014a; Enami et al., 2007).

The purpose of this study is to illustrate the potential of using elastic thermobarometry in combination with structural

and microstructural observations, to better understand the P-T-deformation (D) conditions of retrograde mineral growth in

subduction-related HP/LT metamorphic rocks. We focus on a subduction complex exposed on Syros Island, Cyclades, Greece, 45

where despite several decades of petrological study, the early exhumation history remains enigmatic. We combine the recently

tested quartz-in-epidote (qtz-in-ep) barometer (Cisneros et al., 2020), quartz-in-garnet (qtz-in-grt) barometry (e.g., Ashley et

al., 2014; Bonazzi et al., 2019; Thomas and Spear, 2018), and oxygen isotope thermometry (e.g., Javoy, 1977; Urey, 1947), to

constrain metamorphic growth pressures and temperatures near peak subduction depths and during early exhumation. The

results demonstrate that combining qtz-in-ep barometry with careful structural and microstructural observations allows us to 50

delineate a retrograde P-T-D path that is contextually constrained, and is more robust than what is commonly possible with

conventional thermobarometry.

2. Geologic Setting

Syros Island in the Cyclades of Greece consists of metamorphosed tectonic slices of oceanic and continental affinity

that belong to the Cycladic Blueschist Unit (CBU), structurally below the Pelagonian Upper Unit (Fig. 1). CBU rocks on Syros 55

record Eocene subduction (~52 Ma) to peak blueschist-eclogite facies conditions (Lagos et al., 2007), followed by exhumation

during Oligo-Miocene (~25 Ma) back-arc extension (e.g., Jolivet and Brun, 2010; Ring et al., 2010). A retrograde regional

metamorphic event occurred between 25-18 Ma and caused greenschist- to amphibolite facies metamorphism in the Cycladic

islands, but was most pervasive in the footwall adjacent to the large-scale extensional North and West Cycladic Detachment

Systems (e.g., Bröcker et al., 1993; Bröcker and Franz, 2006; Gautier et al., 1993; Grasemann et al., 2012; Jolivet et al., 2010; 60

Pe-Piper and Piper, 2002; Schneider et al., 2018). Despite these documented metamorphic events, the exhumation history of


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the CBU between ~52 and ~25 Ma remains enigmatic and poorly constrained; yet, this period spans exhumation of the CBU

from maximum subduction to middle crust pressures (~0.3 - 0.7 GPa).

In this work focus on rocks within the CBU, which consist of intercalated metavolcanic and metasedimentary rocks,

metabasites, and serpentinites (e.g., Keiter et al., 2011). Conventional thermobarometry suggests that the CBU on Syros 65

reached peak P-T conditions of ~1.5 GPa and ~500 °C (Ridley, 1984). Trotet et al. (2001a) and Laurent et al. (2018) suggest

higher peak P-T conditions of ~2.0 - 2.4 GPa and ~500 - 550 °C; however, multi-mineral phase equilibria of marbles

(Schumacher et al., 2008) and elastic thermobarometry of metabasites from Kini beach (Behr et al., 2018) support the original

P-T estimates of ~1.5 GPa and 500 °C. Published exhumation P-T paths for the CBU on Syros are also highly variable, ranging

from cooling during decompression, near-isothermal decompression, to reheating during decompression (Laurent et al., 2018; 70

Schumacher et al., 2008; Skelton et al., 2018; Trotet et al., 2001a). The range of previous P-T conditions reflects the lack of

comprehensive studies that combine structural geology, petrology, and thermobarometry across the CBU. Because of these

conflicting P-T paths, several models have been proposed to explain the exhumation history of the CBU, including coaxial

vertical thinning (Rosenbaum et al., 2002), extrusion wedge tectonics (Keiter et al., 2011; Ring et al., 2020), complex thrusting

and extension (Lister and Forster, 2016; Trotet et al., 2001a, 2001b), and subduction channel exhumation (Laurent et al., 2016). 75

3. Field and Microstructural Observations

We studied four localities on Syros (Kalamisia, Delfini, Lotos, Megas Gialos; Fig. 1). Each locality exhibits multiple

stages of mineral growth, and the same deformation and P-T progression. Kalamisia records blueschist facies metamorphism,

and Delfini, Lotos, and Megas Gialos record blueschist-greenschist facies metamorphism. Localities of collected samples and

their associated mineralogy are provided in the supplementary material (Table S1). 80

3.1 Kalamisia

Mafic rocks from Kalamisia preserve retrograde blueschist facies metamorphism (Fig. 1). They exhibit an early

foliation (Ss) characterized by relict blueschist and eclogite facies minerals. The early Ss fabric is re-folded by upright folds

(Ft1) with steeply dipping axial planes, NE-SW-oriented fold hinge lines, and NE-SW-oriented stretching lineations primarily

defined by white mica, glaucophane, and epidote; this indicates folding under blueschist facies conditions (Dt1). 85

Garnets in Kalamisia mafic samples occur as ~1 - 4 mm porphyroblasts (KCS70A, Supplementary Fig. S1a), lack a

well-defined internal foliation, and the Ss foliation deflect around garnets. Glaucophane typically grows within pressure

shadows and brittle fractures of garnet, and omphacite displays breakdown and alteration to glaucophane; this indicates

retrograde glaucophane growth. Glaucophane inclusions within epidote are commonly oriented parallel to Ss, and no omphacite

is observed as inclusions within epidote; these observations support epidote (ep1) growth during retrograde metamorphism. 90


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3.2 Delfini Beach

Metasedimentary rocks at Delfini Beach show retrogression from eclogite- and blueschist- to greenschist facies (Fig.

1). The rocks at Delfini exhibit an early foliation (also considered Ss) characterized by relict blueschist and eclogite facies

minerals (garnet porphyroblasts, and foliation-parallel white mica, blue amphibole, and epidote) aligned in tight isoclinal folds

(Fs) with shallow axial planes. This early fabric was locally retrogressed and re-folded by upright folds (considered Ft2) with 95

steeply dipping axial planes, E-W-oriented fold hinge lines, and E-W-oriented stretching lineations primarily defined by white

mica, chlorite, and actinolite (considered Dt2, Fig. 2a,b); this indicates folding under greenschist facies conditions. Dt2 folding

was associated with boudinage of earlier-generation epidote parallel to the fold hinge lines, and simultaneous precipitation of

new coarse-grained epidote (ep2), along with quartz, calcite and iron oxides in boudin necks (Fig. 3). In some areas of tight

Dt2 folding, a new generation of fine-grained epidote (also interpreted as ep2) grows within a newly developed crenulation 100

cleavage (St, Fig. 2c,d,e).

Garnets in Delfini metasedimentary samples occur as ~1 - 4 mm, partially chloritized porphyroblasts (KCS34, Fig.

2c), and as <1 mm garnets that are commonly found as inclusions within epidote (KCS1621, Supplementary Fig. S1b).

Foliation parallel epidotes (ep1) found within early blueschist-greenschist facies outcrops (KCS1621) range in size from ~0.5

– 5 mm (b-axis length), are strongly poikiloblastic, lack late greenschist facies inclusions such as chlorite, and commonly 105

contain an internal foliation that is oblique to the external matrix Ss foliation (Fig. 2f,g; Supplementary Fig. S1b). Late epidote

(ep2) crystals are found within sample KCS34 from the core of an upright fold (Ft2). During upright folding, a predominant

portion of the rock is recrystallized to late-stage greenschist facies minerals, and contains new epidote (ep2) that is oriented

parallel to the St2 crenulation cleavage. Ep2 crystals range from ~50 - 300 µm along the b-axis (Fig. 2c,d,e), tend to be euhedral

(Fig. 2d,e), sometimes contain titanite inclusions (Fig. 2d), and show textural equilibrium with white mica and titanite that also 110

formed in the St2 cleavage (Fig. 2d,e). Ep2 crystals are not poikiloblastic and rarely preserve quartz inclusions, thus only a few

analyses were possible.

3.3 Lotos Beach

The rocks from Lotos Beach exhibit the same structural and petrological progression as those from Delfini (Fig. 1),

showing retrogression from eclogite- and blueschist- to greenschist facies. An early Ss foliation was locally retrogressed and 115

re-folded by upright Ft2 folds with steeply dipping axial planes, E-W-oriented fold hinge lines, and E-W-oriented stretching

lineations primarily defined by white mica, chlorite, and actinolite (Dt2). Dt2 folding was associated with boudinage of earlier-

generation epidote parallel to the fold hinge lines, and simultaneous precipitation of new coarse-grained epidote (ep2), along

with quartz, calcite and iron oxides in boudin necks (Fig. 3).

Garnets in Lotos samples occur as ~1 - 3 mm chloritized porphyroblasts (e.g., KCS3), that deflect the external Ss 120

foliation (KCS3). Foliation parallel epidotes (ep1) found within early blueschist-greenschist facies outcrops (SY1402, SY1405,

KCS2, KCS3) range in size from ~0.5 – 5 mm (b-axis length), are strongly poikiloblastic, and commonly contain an internal


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foliation that is oblique to the external matrix Ss foliation (Supplementary Fig. S1c). Boudinage of ep1 parallel to stretching

lineations is common in thin sections (Supplementary Fig. S1c).

3.4 Megas Gialos 125

The rocks from Megas Gialos exhibit the same structural and petrological progression as those from Lotos and Delfini

Beaches (Fig. 1). Rocks from Megas Gialos show retrogression from eclogite- and blueschist- to greenschist facies. An early

Ss foliation was locally retrogressed and stretching lineations primarily defined by white mica, chlorite, and actinolite are E-

W-oriented.

No garnets were found within the analyzed sample from Megas Gialos. Foliation parallel epidotes (ep1) found within 130

early blueschist-greenschist facies outcrops range in size from ~0.5 – 3 mm (b-axis length), are strongly poikiloblastic, and

commonly contain an internal foliation that is oblique to the external matrix Ss foliation (Supplementary Fig. S1d). Boudinage

of ep1 parallel to stretching lineations is common in thin sections (Supplementary Fig. S1d).

4. Methods

We determined P-T conditions using elastic thermobarometry and oxygen isotope thermometry. Raman spectroscopy 135

was used to measure Raman shifts of strained quartz inclusions entrapped within epidote or garnet, and a laser fluorination

line and a GasBench II coupled to a gas source mass spectrometer was used to measure oxygen isotope ratios of quartz and

calcite separates, respectively.

4.1 Raman Spectroscopy measurements

Our Raman spectroscopy measurements are taken from ~30 µm, ~80 µm, and ~150 µm thin and thick sections, that 140

consist of sections cut perpendicular to foliation (Ss) and parallel to stretching lineations (e.g., KCS1621), and perpendicular

to the Ft2 fold axial plane (KCS34). Quartz inclusions were measured from multiple epidotes and garnets within individual

sections (Supplementary Table S3). Measured quartz inclusions were small in diameter relative to the host, and were two-to-

three-times the inclusion radial distance from other inclusions, fractures, and the host exterior to avoid overpressures or stress

relaxation (Fig. 4a,b; Campomenosi et al., 2018; Zhong et al., 2020). No geometric corrections were applied (Mazzucchelli et 145

al., 2018).

Raman spectroscopy measurements of quartz inclusions within garnet and epidote were carried-out at Virginia Tech

(VT) and ETH Zürich (ETHZ) by using JY Horiba LabRam HR800 and DILOR Labram Raman systems, respectively.

Analyses at VT used an 1800 grooves mm-1 grating, 100x objective with a 0.9 numerical aperture (NA), 400 µm confocal

aperture, and a 150 µm slit width. Raman spectra were centered at ~360 cm-1. We used a 514.57 nm wavelength Ar laser, and 150

removed the laser interference filter for all analyses to apply a linear drift correction dependent on the position of the 116.04

cm-1, 266.29 cm-1, and 520.30 cm-1 Ar plasma lines (Fig. DR4). Measurements at ETHZ used a 532 nm laser, an 1800 grooves


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mm-1 grating, a 100x objective with a 0.9 NA, a 200 µm confocal aperture, and a 300 µm slit width. Raman spectra were

centered at ~ 850 cm-1.

All Raman spectra was reduced with a Bose-Einstein temperature-dependent population factor (Kuzmany, 2009). All 155

Raman bands were fit by using PeakFit v4.12 from SYSTAT Software Inc. A Gaussian model was used to fit Ar plasma lines

(only VT analyses), and a Voigt model was used to fit the quartz 128 cm-1, 206 cm-1, and 464 cm-1 bands, epidote bands, and

garnet bands. Raman bands of quartz, epidote, and garnet, and Ar plasma lines were fit simultaneously, and a linear background

subtraction was applied during peak fitting. Baseline-to-baseline deconvolution of quartz and garnet bands was simple and

generally required fitting quartz bands and a few shoulder garnet bands. Deconvolution of quartz and epidote bands required 160

more complicated deconvolution; we followed a fitting approach similar to that described by Cisneros et al. (2020).

4.2 Inclusion and entrapment pressure calculations

The fully encapsulated inclusions preserve strain that causes the Raman active vibrational modes of inclusions to be

shifted to higher or lower wavenumbers relative to minerals that are unstrained (fully exposed). We calculated the Raman

shift(s) of inclusions (ωinc) relative to Raman shift(s) of an unencapsulated Herkimer quartz standard (ωref) at ambient 165

conditions (Δω = ωinc – ωref) (Fig. 4). For VT analyses, ωinc was measured relative to a Herkimer quartz standard that was

analyzed 5 times prior to same day analyses. A drift correction was applied to ωinc by monitoring the position of Ar plasma

lines (Supplementary Tables S2; S3). For ETHZ analyses, a Herkimer quartz standard was analyzed 3 times prior to and after

quartz inclusion analyses. A time-dependent linear drift correction was applied to ωinc based on the drift shown by Herkimer

quartz analyses that bracketed inclusion analyses (< 0.2 cm-1). 170

We calculated residual inclusion pressures (Pinc) by using hydrostatic calibrations and by accounting for quartz

anisotropy. To calculate a Pinc from individual quartz Raman bands, we used pressure-dependent Raman shift(s) (P-Δω) of the

quartz 128 cm-1, 206 cm-1, and 464 cm-1 bands, that have been experimentally calibrated under hydrostatic stress conditions

by using diamond anvil cell experiments (Schmidt and Ziemann, 2000). To account for quartz anisotropy, we calculated Pinc

from strains. Calculating quartz strains requires that the Raman shift of at least 2 quartz vibrational modes can be measured. 175

When we were able to measure the quartz 128, 206 and 464 cm-1 band positions of inclusions, we calculated strains from the

Δω of 3 bands. If only two bands were measured, we calculated strains from the Δω of 2 bands (Supplementary Table S3). For

the remaining analyses with low 128 and 206 cm-1 intensities, we report Pinc calculated from the 464 cm-1 band hydrostatic P-

Δω relationship (Supplementary Table S3). Strains were determined from the Δω of the 128 cm-1, 206 cm-1, and 464 cm-1

quartz bands by using Strainman (Angel et al., 2019; Murri et al., 2018, 2019), wherein a weighted fit was applied based on 180

the Δω error associated with each quartz Raman band. Calculated strains were converted to a mean stress [Pinc = (2σ1 + σ3)/3]

using the matrix relationship σi = cijεj, where σi, cij, and εj, are the stress, elastic modulus, and strain matrices, respectively. We

used the α-quartz trigonal symmetry constraints of Nye (1985) and quartz elastic constants of Wang et al. (2015).

We assumed constant mineral compositions for all modeling (epidote: Xep = 0.5 and Xcz = 0.5; garnet: XAlm = 0.7,

XGr = 0.2, and XPy = 0.1). Garnet compositions have a negligible effect on entrapment pressures (Ptrap) because the 185


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thermodynamic and physical properties of garnet end-members are similar (e.g., Supplementary Table S8); however, epidote

composition has a greater effect on Ptrap (Cisneros et al., 2020). To account for epidote and garnet solid solutions, we

implemented linear mixing of shear moduli and molar volumes (V). Ideal mixing of molar volumes has been shown to be an

appropriate approximation for epidote-clinozoisite solid solutions (Cisneros et al., 2020; Franz and Liebscher, 2004). Garnet

molar volumes were modeled using the thermodynamic properties of Holland and Powell (2011) (almandine and pyrope) and 190

Milani et al. (2017) (grossular), and a Tait Equation of State (EoS) with a thermal pressure term. We used the shear moduli of

Wang and Ji (2001) (almandine and pyrope) and Isaak et al. (1992) (grossular). Epidote molar volumes were modeled using

the thermodynamic properties and shear moduli given by Cisneros et al. (2020), and a Tait EoS and thermal pressure term.

Epidote and clinozoisite regressions are based on the P-V-T data of Gatta et al. (2011) (Xep = 0.74), and T-V and P-V data of

Pawley et al. (1996) (Xep = 0) and Qin et al. (2016) (Xep = 0.39), respectively. Clinozoisite and epidote have similar thermal 195

expansivities but differing bulk moduli (Supplementary Table S4). To account for the composition of epidotes used in P-V-T

experiments, we normalized the composition of our unknown epidotes across the compositional range of P-V experimental

epidotes, i.e., the molar volume of our unknown epidote (Xep = 0.5) is estimated as 31 % (Xep = 0.74) and 69 % (Xep = 0.39)

of each experimental epidote. Quartz molar volumes were modeled using the thermodynamic properties and approach of Angel

et al. (2017a). Entrapment pressures were calculated from residual quartz Pinc by using the Angel et al. (2017b) 1D elastic 200

model equation, and a MATLAB program available in Cisneros and Befus (2020) that implements mixing of shear moduli and

molar volumes. A comparison of entrapment pressures calculated from the Cisneros and Befus (2020) MATLAB program and

EoSFit-Pinc (Angel et al., 2017b) is given in Supplementary Table S4; entrapment pressure calculations of mineral end-

members accounts for the reproducibility of molar volume and elastic modeling calculations.

4.3 Stable isotope measurements 205

Samples were measured by using a ThermoElectron MAT 253 isotope ratio mass spectrometer (IRMS) at the

University of Texas at Austin. Quartz δ18O values were measured by laser fluorination (Sharp, 1990), and ~2.0 mg of quartz

were used in each analysis. Quartz from samples SY1613, SY1617, and SY1623 was duplicated to determine isotopic

homogeneity and reproducibility. An internal quartz Lausanne-1 standard (δ18O = +18.1‰) was analyzed with all samples to

evaluate precision and accuracy. All δ18O values are reported relative to standard mean ocean water (SMOW), where the δ18O 210

value of NBS-28 is +9.65‰. Measurement precision based on the long-term reproducibility of standards is ± 0.1 ‰ (1 σ).

Precision of Lausanne-1 on the day of analysis was ± 0.3 ‰ (1 σ), whereas samples reproduced with a precision of ± 0.1 ‰

(1 σ) or better (Supplementary Table S5). Calcite δ18O values were measured on a Thermo Gasbench II coupled to a

ThermoElectron 253 mass spectrometer. Each analysis used 0.25 - 0.5 mg of calcite that was loaded into Exetainer vials,

flushed with ultra-high purity helium, and reacted with 103 % phosphoric acid at 50 °C for ~2 hours. Headspace CO2 was then 215

transferred to the mass spectrometer. Samples were calibrated to an in-house standard, NBS-18, and NBS-19. Measurement

precision is ± 0.04 ‰ (1 σ) based on the long-term reproducibility of standards.


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4.4 Stable isotope temperature calculations

Temperatures derived from stable isotope measurements were calculated by using the Sharp and Kirschner (1994)

quartz-calcite oxygen isotope fractionation calibration (A = 0.87 ± 0.06; equation A1; Supplementary Table S5). Isotopic 220

equilibrium was assumed for all samples. Several observations support that this assumption is appropriate: 1) duplicate δ18O

analysis of quartz and calcite grains give the same isotopic value, suggesting grain isotopic homogeneity, 2) the stage of

deformation that these mineral pairs are related to is not affected by further deformation in either outcrop or thin section, and

3) all quartz-calcite pairs suggest a similar temperature of isotopic equilibration.

Temperature errors from quartz-calcite oxygen isotope measurements were calculated through the square-root of the 225

summed quadratures of all sources of uncertainty (equations A2, A3). These uncertainties included δ18O value errors of quartz

and calcite of ± 0.1 ‰ (1 σ) and ± 0.04 ‰ (1 σ), respectively, and errors associated with the Sharp and Kirschner (1994)

quartz-calcite oxygen isotope fractionation calibration (A parameter).

4.5 Electron probe measurements

Electron probe analyses were carried-out at ETHZ using a JEOL JXA-8230 Electron Probe Microanalyzer (EPMA). 230

The EPMA is equipped with five wavelength-dispersive spectrometers. Epidote and omphacite were analyzed for Si Al, Na,

Mg, Ca, Cr K, Ti, Fe, and Mn on TAP, TAPH, PETJ, PETL, and LIFH crystals. Beam parameters included a 20 nA beam

current, 10 um beam size, and a 15 keV accelerating voltage. All elements were measured for 30 s on peak and a mean atomic

number background correction was applied. Primary standards used included: albite, anorthite, synthetic forsterite, chromite,

microcline, synthetic rutile, synthetic fayalite, and synthetic pyrolusite. 235

5. Results

Determined pressures were categorized into three groups according to outcrop and microstructural context (Fig. 5;

Fig. 6; Supplementary Table S3): garnet growth near peak metamorphic conditions (Group 1), growth of foliation-parallel

epidote during blueschist-greenschist facies metamorphism (ep1, Group 2), and late-stage epidote growth in the new

crenulation (St2) associated with Ft2 folds during greenschist facies metamorphism (ep2, Group 3). New ep2 growth is also 240

supported by the mineral chemistry of different epidote generations within the St2 crenulation. Epidotes show a progressive

chemical evolution that is recorded by an early generation epidote inclusion in titanite that occurs parallel to St2 (Xep ≅ 0.1),

the ep2 core (Xep ≅ 0.5), and the ep2 rim (Xep ≅ 0.8) (Fig. 2g; Supplementary Table S6).

The entrapment temperature (Ttrap) of quartz inclusions in garnet (garnet growth temperature) is estimated as 500 -

550 °C; this is based on good agreement between previous studies on the maximum temperature reached by CBU rocks from 245

Syros (e.g., Laurent et al., 2018; Ridley, 1984; Schumacher et al., 2008; Skelton et al., 2018; Trotet et al., 2001a). Ttrap for the

ep2 population (Group 3) is deduced from oxygen isotope thermometry of quartz-calcite boudin-neck precipitates. The mean

temperature from quartz-calcite pairs from boundin necks is 411 ± 23 °C (n = 4, Supplementary Table S5). Ttrap for the ep1


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population (Group 2) is estimated as being intermediate between garnet and ep2 growth (~400 - 500 °C). As shown by qtz-in-

ep isomekes (constant Pinc lines along which fractional volume changes of an inclusion and host are equal), the assumed Ttrap 250

has a minimal effect on Ptrap (Fig. 6b; Cisneros et al., 2020).

5.1 Kalamisia

Group 1 quartz-inclusions-in-garnet record a mean Pinc of 600 ± 78 MPa (Fig 4; Supplementary Table S3). This

corresponds to an entrapment pressure (Ptrap) of 1.43 - 1.49 ± 0.14 GPa (n = 5), at an estimated Ttrap between 500 - 550 °C (Fig.

6b, Supplementary Table S3). Group 2 quartz-inclusions-in-ep1 record a mean Pinc of 544 ± 57 MPa, corresponding to a Ptrap 255

of 1.43 ± 0.12 (n = 6) at an estimated Ttrap of 450 °C. No Group 3 epidotes are found within our analyzed section from

Kalamisia.

5.2 Delfini

Group 1 records a mean Pinc of 739 ± 49 MPa (Fig 4; Supplementary Table S3). This corresponds to a Ptrap of 1.68 -

1.74 ± 0.09 GPa (n = 20), at an estimated Ttrap between 500 - 550 °C (Fig. 6b, Supplementary Table S3). Group 2 records a 260

mean Pinc of 518 ± 52 MPa, corresponding to a Ptrap of 1.38 ± 0.11 (n = 5) at an estimated Ttrap of 450 °C. Group 3 records a

mean Pinc of 343 ± 23 MPa, corresponding to a Ptrap of 0.98 ± 0.05 GPa (n = 3) at 411 °C (Supplementary Table S3).

5.3 Lotos

Group 1 records a mean Pinc of 751 ± 76 MPa (Fig 4; Supplementary Table S3). This corresponds to a Ptrap of 1.70 -

1.76 ± 0.14 GPa (n = 2), at an estimated Ttrap between 500 - 550 °C (Fig. 6b; Supplementary Table S3). Group 2 records a 265

mean Pinc of 531 ± 78 MPa, corresponding to a Ptrap of 1.41 ± 0.17 (n = 15) at an estimated Ttrap of 450 °C. No Group 3 epidotes

were analyzed from Lotos.

5.4 Megas Gialos

Group 2 records an average Pinc of 494 ± 29 MPa (Fig. 6), corresponding to a Ptrap of 1.33 ± 0.03 (n = 6) at an estimated

Ttrap of 450 °C (Fig. 6b; Supplementary Table S3). No Group 1 garnets or Group 3 epidotes were analyzed from Megas Gialos. 270

6. Discussion

6.1 Elastic thermobarometry pressure groups

Group 1 garnets either lack an internal foliation or contain a weak foliation that is defined by inclusions oblique to

the Ss fabric, which indicates a previous stage of deformation (Fig. 2c; Supplementary Fig. S1a,b). Garnets record similar

pressures, regardless of the location of quartz inclusions (Supplementary Table S3). Omphacite inclusions within different 275


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garnet zones also show no difference in composition, which is consistent with qtz-in-grt barometry results (Delfini: KCS1621,

Supplementary Table S5). Group 2 epidotes (ep1) overgrow garnets, are aligned parallel to the Ss foliation but sometimes

preserve an internal foliation that is oblique to Ss, and lack late greenschist facies inclusions (Fig. 2f,g; Supplementary Fig.

S1). Group 3 epidotes (ep2, KCS34, Fig. 2c,d,e) are short in length, are aligned parallel to a late St2 crenulation, contain

minimal quartz inclusions, and only record Group 3 pressures, independent of the position of quartz inclusions within epidotes. 280

Based on these observations, the Group 1 Ptrap estimates from the qtz-in-grt barometer record high-P conditions on

Syros associated with prograde garnet growth, and the Group 2 and 3 Ptrap estimates from the qtz-in-ep barometer record epidote

growth during early blueschist-greenschist facies retrogression (ep1, Dt1) and subsequent Dt2 deformation (ep2), respectively.

We interpret the low-P epidote group (Group 3) to be associated with Dt2 folding, and best recorded in areas that experienced

late greenschist facies mineral growth due to enhanced deformation and/or fluid influx during this stage of deformation (e.g., 285

core of Ft2 fold).

6.2 Comparison of peak pressure constraints for the CBU on Syros and Sifnos

Based on qtz-in-grt measurements (Group 1), our Ptrap calculations suggest maximum P conditions of ~1.6 - 1.8 GPa

were reached by the CBU on Syros. Garnets from metasedimentary and metavolcanic rocks record the statistically highest Ptrap

(~1.5 - 1.8 GPa), whereas garnets from metamafic rocks (Kalamisia) record the lowest Ptrap (~1.3 - 1.6 GPa) (Fig. 6b). We 290

present a compilation of previous P-T constraints on CBU rocks from Syros and Sifnos, Greece, and discuss how our Ptrap

constraints compare with previous studies.

Elastic thermobarometry, mineral stability constraints, and multi-phase equilibrium modeling results from Sifnos

CBU rocks suggest maximum P conditions of ~1.8 ± 0.1 GPa (Ashley et al., 2014), ~1.4 ± 0.2 GPa (Matthews and Schliestedt,

1984), and ~2.0 - 2.2 GPa (Dragovic et al., 2012, 2015; Groppo et al., 2009; Trotet et al., 2001a), respectively. The results of 295

Ashley et al. (2014) are commonly cited as evidence that the CBU reached high pressure conditions (≥ 2.0 GPa, from elastic

thermobarometry); however, their Ptrap calculations were carried out by using fits to quartz molar volume (P-T-V) data that

have recently been re-evaluated (Angel et al., 2017a) . Improved fits to quartz molar volume experiments “soften” quartz, and

remodeling Pinc values from Ashley et al. (2014) reduces maximum mean Ptrap conditions to ~1.6 ± 0.1 GPa (Fig. 6a,

Supplementary Table S7). 300

Elastic thermobarometry, mineral stability constraints, glaucophane-bearing marble mineral equilibria, and multi-

phase equilibria modeling results from Syros CBU rocks suggest peak pressure conditions of ~1.5 ± 0.1 GPa (Behr et al.,

2018), ~1.4 - 1.9 GPa (Ridley, 1984), ~1.5 GPa (Schumacher et al., 2008), and ~1.9 - 2.4 GPa (Laurent et al., 2018; Skelton

et al., 2018; Trotet et al., 2001a), respectively. Elastic thermobarometry results from Syros, Greece were reduced using the

approach outlined in (Ashley et al., 2016), wherein a correction to Ptrap is applied based on the assumed Ttrap. Recent studies 305

suggest that not using a temperature-dependent Ptrap correction produces suitable results that accurately reproduce experimental

conditions of quartz entrapment by garnet (Bonazzi et al., 2019; Thomas and Spear, 2018). Recalculation of the Behr et al.

(2018) Pinc data (no temperature-dependent Ptrap correction) results in a mean Ptrap of ~1.7 ± 0.1 GPa (Fig. 6a, Supplementary


11

Table S8). The re-evaluation of data from Ashley et al. (2014) and Behr et al. (2018) suggests that our results are in good

agreement with previous elastic thermobarometry constraints, and that to date, no qtz-in-grt elastic thermobarometry results 310

suggest pressures ≥ 2.0 GPa.

Different methodologies applied to CBU rocks from Syros have resulted in a wide range of maximum P estimates.

Schumacher et al. (2008) used mineral-equilibria modeling of glaucophane-bearing marbles to place constraints on maximum

P-T conditions. Maximum P-T conditions are constrained by the presence of glaucophane + CaCO3 + dolomite + quartz, which

suggests that the marbles exceeded the albite/Na-pyroxene + dolomite + quartz → glaucophane + CaCO3 reaction, but did not 315

cross the dolomite + quartz → tremolite + CaCO3 or the glaucophane + aragonite-out reactions. The mineral reaction

constraints suggest maximum P-T conditions of ~ 1.5 - 1.6 GPa and 500 °C for the CBU marbles. Ridley (1984) used the

stability of paragonite and lack of kyanite to deduce max P constraints of ~1.4 -1.9 GPa. Trotet et al. (2001b, 2001a), Laurent

et al. (2018), and Skelton et al. (2018) employed thermodynamic phase-equilibria modeling and supplementary methods to

constrain P-T conditions for CBU rocks from Syros. Skelton et al. (2018) used the Powell and Holland (1994) Thermocalc 320

database, Trotet et al. (2001b, 2001a) used the Berman (1991) thermodynamic database and the TWEEQC approach, and

Laurent et al. (2018) used empirical thermobarometry, GrtMod (Lanari et al., 2017), and isochemical phase diagrams. Trotet

et al. (2001b, 2001a), Laurent et al. (2018), and Skelton et al. (2018) found high-P conditions for the CBU (≥ 1.9 GPa), and

results from Laurent et al. (2018) suggest some rocks reached conditions as high as 2.4 GPa. Results from Laurent et al. (2018)

suggest most garnet growth occurred at ~1.8 GPa and 500 °C; however, some garnet modeling results suggest that garnet rims 325

grew at ~2.4 GPa and 500 - 550 °C, albeit errors are increasingly large for these results (± 0.4 - 0.9 GPa).

Some GrtMod results suggest prograde core and rim garnet growth at ~1.8 GPa and 475 °C, and ~2.4 GPa and 475

°C, respectively (sample SY1418 from; Laurent et al., 2018). These results would indicate that the garnets grew under

isothermal conditions during prograde subduction, a result that we find difficult to reconcile within a subduction zone

geothermal gradient. Garnet results from another sample (SY1401) suggest core and rim garnet growth at ~1.8 GPa and 475 330

°C, and ~2.4 GPa and 550 °C, respectively. Sample SY1401 is collected from the same locality as ours (Kalamisia), but our

qtz-in-grt results from this study suggest that garnets from this outcrop record the statistically lowest Ptrap. It is possible,

however, that we did not sample the same rocks as Laurent et al. (2018), or that we have not found or analyzed garnets that

record high pressures.

Previous studies have also suggested that pressures ≥ 2.0 GPa are unreasonable for Syros because paragonite is 335

abundant in CBU rocks, but kyanite has not been reported. This suggests that CBU rocks did not cross the reaction paragonite

→ jadeite50 + kyanite + H2O (~1.9 - 2.0 GPa); however, we recognize that the occurrence of kyanite may require high

Al2O3:SiO2 ratios for metabasites (e.g., Liati and Seidel, 1996), and that the pressure of this reaction is compositionally

dependent. Pseudosections of eclogite CBU rocks show that kyanite would not be found in these bulk compositions below

~2.3 GPa (Skelton et al., 2018). It is possible that the high-P conditions found in previous studies may be real, but may only 340

be recorded locally within some eclogite blocks.


12

In general, phase stability relationships (e.g., Matthews and Schliestedt, 1984; Ridley, 1984; Schumacher et al., 2008)

and qtz-in-grt barometry results are in good agreement, but do not agree with high-pressure results (≥ 1.9 GPa) deduced from

thermodynamic modeling using approaches such as GrtMod and TWEEQC. The difference between our results and those of

previous studies is important to reconcile, because the maximum P conditions reached by the CBU has considerable 345

implications for the internal architecture of the CBU, its geodynamic evolution, and the mechanisms that can accommodate

exhumation mechanisms of high-P subduction zone rocks from Syros. A comparison of qtz-in-grt barometry with

thermodynamic modeling results from samples that record high pressures would be appropriate for further testing differences

between the two techniques.

6.3 Comparison of exhumation P-T conditions 350

Previous studies have presented varying P-T paths and associated exhumation histories for Syros CBU rocks (Fig.

6a; Laurent et al., 2018; Schumacher et al., 2008; Skelton et al., 2018; Trotet et al., 2001a). We present a compilation of

previous P-T constraints and interpretations and discuss how our results compare with previous studies.

Schumacher et al. (2008) do not provide quantitative constraints for the retrograde P-T path (schematic), and samples

do not have structural context; however, the authors suggest that a “cold” P-T path during exhumation is required for Syros 355

CBU rocks based on the occurrence of lawsonite + epidote assemblages across Syros, and the P-T path required to avoid

crossing the lawsonite→ kyanite + zoisite reaction (Fig. 6a). The authors suggest that exhumation of CBU packages occurred

shortly after juxtaposition near peak metamorphic conditions.

Both Trotet et al. (2001a, 2001b) and Laurent et al. (2018) constrain high-P conditions for the CBU (> 2.0 GPa),

however, their proposed exhumation histories differ. Trotet et al. (2001b) suggested that CBU eclogites, blueschists and 360

greenschists underwent different T-t histories during exhumation and were juxtaposed late along ductile shear zones. Laurent

et al. (2018) suggested that the entire CBU reached peak metamorphic conditions of ~2.2 GPa, and that units that preserved

blueschist facies assemblages underwent cooling during decompression, whereas rocks of southern Syros from lower structural

levels experienced isobaric heating (~550 °C) at mid-crustal depths (~1.0 GPa) followed by subsequent cooling. Laurent et al.

(2018) interpreted reheating to indicate that CBU rocks on Syros reached high-P conditions, and then rapidly transitioned from 365

a forearc to back-arc setting, thus experiencing a period of increasing temperatures.

Skelton et al. (2018) also estimated peak and exhumation P-T conditions of rocks from Fabrikas (southern Syros),

and interpreted exhumation of the CBU within an extrusion wedge (Ring et al., 2020). The authors constrained maximum P-

T conditions of ~1.9 GPa and 525 °C, and retrograde conditions of ~1.4 GPa and 500 °C (blueschist facies) and ~0.3 GPa and

450 °C (greenschist facies) based on Thermocalc end-member activity modeling (Powell and Holland, 1994). Retrograde 370

blueschist conditions (inferred from garnet growth) are similar between their estimates and ours, but greenschist facies

conditions vastly differ. However, Skelton et al. (2018) focused on greenschist facies outcrops wherein metamorphism

occurred locally over short length scales (e.g. ~10 - 100 m), adjacent to late-stage brittle normal faults. We interpret our Dt2

stage of greenschist facies metamorphism to pre-date late-stage normal faulting that has been attributed to Neogene block


13

rotations (Cooperdock and Stockli, 2016) or possible coeval granitoid magmatism during Miocene back-arc extension (Keiter 375

et al., 2011).

Our results show that rocks from Kalamisia, Delfini, Lotos, and Megas Gialos, reached peak P-T conditions and

underwent cooling during retrograde blueschist and greenschist facies metamorphism (Fig. 6b). Peak P-T conditions of the

CBU are ~1.6 - 1.8 GPa and 500 - 550 °C (Group 1 qtz-in-grt Ptrap estimates), indicating a subduction zone geothermal gradient

of ~9 - 10 °C km-1 at ~55 - 60 km (assuming 30 MPa km-1). Group 2 and 3 qtz-in-ep Ptrap estimates indicate geothermal gradients 380

of ~10 °C km-1 and ~12 °C km-1 at ~47 and 33 km depths, respectively (Fig. 6b), demonstrating a similar P-T trajectory during

exhumation. Our P-T constraints are inconsistent with reheating to ~550 °C and 1.0 GPa, wherein amphibolite facies

mineralogy may be stable. Our samples and the sample from which Laurent et al. (2018) determined reheating (SY1407),

preserve no mineralogical evidence for having reached epidote-amphibolite facies (Fig. 6a; e.g., pargasite/hornblende,

biotite/muscovite). Instead, the matrix mineralogy of sample SY1407 (glaucophane, phengite, rutile) suggests that these rocks 385

formed under a cold geothermal gradient, rather than in a back-arc setting with an elevated geothermal gradient. Laurent et al.

(2018) suggest that sample SY1407 records albite-epidote-blueschist conditions, a field metamorphic facies that can expand

to higher T conditions; however, a pseudosection created for a similar bulk composition suggests that the determined P-T

constraints (~1.0 GPa and 550 °C) are within epidote-amphibolite facies (Trotet et al., 2001a). Furthermore, results from

sample SY1407 of Laurent et al. (2018) significantly disagree when using local vs. bulk compositions for modeling. Models 390

that use bulk compositions suggest that the core and mantle of the garnet record P-T conditions ~1.8 GPa and 475 °C, whereas

models that use local compositions suggest that the garnets do not record conditions above ~1.0 GPa. Lanari and Engi (2017)

have documented this issue, and describe how implemented compositions can drastically affect calculated P-T conditions; it

remains unclear how much the chosen bulk composition alters the P-T constraints of Laurent et al. (2018) for sample SY1407.

Our results suggest that rocks from different Syros outcrops record similar peak and exhumation P-T conditions, but 395

experienced different extents of deformation and thus recrystallization during exhumation. The similar peak pressures between

different Syros outcrops suggests that these rocks belong to what has been called the “Upper Cycladic Blueschist Nappe” on

Milos Island (as opposed to the “Lower Cycladic Blueschist Nappe), which records peak pressure conditions above ~0.8 GPa

(Grasemann et al., 2018). The observation of similar P-T conditions reached at different locations is inconsistent with results

that suggest individual P-T paths for rocks that preserve different metamorphic facies (Trotet et al., 2001b, 2001a), and 400

different sections of the CBU (Laurent et al., 2018); however, we do not have T constraints for rocks from southern Syros. Our

results are in better agreement with a P-T evolution resembling that of Schumacher et al. (2008), and a geothermal gradient of

~10 – 12 °C km-1 that has also been proposed for CBU rocks from Sifnos, Greece (Schmädicke and Will, 2003).

6.4 Implications for exhumation mechanisms

Our results indicate that the CBU followed a “cooling during decompression” P-T trajectory that required a heat sink 405

at depth to cool rocks during exhumation. Cooling could be achieved under a steady-state subduction zone thermal gradient

with slab-top temperatures similar to those of warm subduction zones, such as in Cascadia (e.g., Syracuse et al., 2010;


14

Walowski et al., 2015). This would suggest that exhumation was achieved parallel to the subducting plate, in a subduction

channel geometry prior to core-complex formation. During this phase of exhumation, CBU rocks remained within a cold

forearc until they reached the mid-crust (~1.0 GPa), and exhibit a progressive change in kinematics, from N-S stretching 410

lineations during subduction (Behr et al., 2018), to lineations that swing towards the E-W during exhumation (c.f., Kotowski

and Behr, 2019; Laurent et al., 2016). The inferred P-T conditions and kinematics of our studied samples are consistent with

Syros recording early deformation and metamorphism within a forearc setting, whereas adjacent Cycladic islands that border

the North and West Cycladic Detachment Systems record late-stage kinematics and greenschist facies metamorphism that

capture the CBU transition to a warmer back-arc setting. 415

7. Conclusions

This work highlights the potential of using elastic thermobarometry in combination with structural (macro and micro)

and petrographic constraints, to better constrain P-T conditions of challenging rock assemblages. Our results allow us to place

robust P-T constraints on distinct textural fabrics that are related to well-constrained outcrop scale structures. In particular, the

work highlights how the qtz-in-ep barometer is well suited for constraining formation conditions of epidote, a common mineral 420

that is found within a large range of geologic settings and P-T conditions. Combining the qtz-in-ep barometer with other elastic

thermobarometers (e.g., qtz-in-grt) allows determination of protracted P-T histories from minerals that record different

geologic stages within single rocks samples.

Our new results show that CBU rocks from Syros, Greece, experienced similar P-T conditions during subduction and

exhumation, inconsistent with results that suggest different P-T histories for CBU rocks for Syros or increasing temperatures 425

during exhumation. Our targeted stages of deformation and metamorphism suggest that CBU rocks from Syros record cooling

during decompression, consistent with exhumation within a subduction channel and early deformation and metamorphism

within a forearc (at least to ~33 km depth), prior to Miocene core-complex formation and transition to a warmer back-arc

setting.

Appendix A: Stable isotope temperature error calculations 430

Temperature errors from oxygen isotope measurements were calculated through the square-root of the summed

quadratures of all sources of uncertainty. These uncertainties included error of δ18O values of quartz (qtz) and calcite (cc) of ±

0.1 ‰ (1 σ) and ± 0.04 ‰ (1 σ), respectively, and errors associated with the Sharp and Kirschner (1994) quartz-calcite oxygen

isotope fractionation calibration (A parameter). Errors from the sum of propagated analytical errors, were propagated through

the empirical calibration of quartz-calcite oxygen isotope fraction that was used for temperature calculations: 435

∆𝑞𝑞𝑞𝑞𝑞𝑞−𝑐𝑐𝑐𝑐=𝐴𝐴 × 106

𝑇𝑇2 𝐴𝐴1


15

where A = 0.87 ± 0.06 (1 σ). The square-root of the summed quadratures is expressed as:

𝜎𝜎𝑇𝑇 = �𝜎𝜎𝐴𝐴2 �𝜕𝜕𝑇𝑇

𝜕𝜕𝐴𝐴�

2

+ 𝜎𝜎∆𝑞𝑞𝑞𝑞𝑞𝑞−𝑐𝑐𝑐𝑐2 �

𝜕𝜕𝑇𝑇

𝜕𝜕∆𝑞𝑞𝑞𝑞𝑞𝑞−𝑐𝑐𝑐𝑐�

2

𝐴𝐴2

440

𝜎𝜎𝑇𝑇 = �𝜎𝜎𝐴𝐴2 �0.5 ∗ 103

√𝐴𝐴 ∗ �∆𝑞𝑞𝑞𝑞𝑞𝑞−𝑐𝑐𝑐𝑐�

2

+ 𝜎𝜎∆𝑞𝑞𝑞𝑞𝑞𝑞−𝑐𝑐𝑐𝑐2 �−0.5 ∗

√𝐴𝐴 ∗ 103

∆𝑞𝑞𝑞𝑞𝑞𝑞−𝑐𝑐𝑐𝑐1.5�2

𝐴𝐴3

Author Contribution

All authors contributed to this manuscript. M. Cisneros developed the epidote barometer, collected the data, and wrote

the manuscript. J. Barnes, W. Behr, A. Kotowski, D. Stockli, and K. Soukis helped with conceiving the project, field work,

and writing. 445

Acknowledgements

We thank N. Raia for field work assistance, J. Allaz for assistance on the microprobe at ETH Zürich, and C. Farley

and R. Bobnar for access to the Raman Spectrometer at Virginia Tech. This work was supported by a GSA Student Research

Grant and a Ford Foundation Fellowship awarded to M.C, an NSF Graduate Research Fellowship awarded to A.K., and NSF 450

Grant (EAR-1725110) awarded to J.B., W.B., and D.S.

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Figure 1. Geologic map of Syros, Greece [modified from Keiter et al. (2011) and Behr et al. (2018)]. Inset map shows Syros relative 640 to the North and West Cycladic, and Naxos-Paros Detachment Systems (NCSD, WCDS, NPDS, modified from Grasemann et al., 2012). Stereonets from each studied outcrop are shown, and arrows indicate the outcrop location.


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Figure 2. Outcrop, micrograph, and electron images showing stages of retrograde deformation present in southern Delfini. a) 645 Upright folds (Ft2) that refold the primary Ss foliation. b): Core of Ft2 folds (below Fig. 2a, KCS34). c): Plane light image of sample KCS34; sample cut perpendicular to the Ft2 fold axial plane. Epidotes (ep2) from the upright fold exhibit recrystallization as indicated by alignment with a late St2 crenulation, and a reduction in inclusions and grain size. d) Ep2 with late titanite (ttn) inclusions. Ep2 is parallel to white mica (wm) that defines St2 (KCS34). e) Ep2 in textural equilibrium with ttn (KCS34). f) Ep1 parallel to Ss, with garnet (grt) and quartz (qtz) inclusions that do not define an internal foliation (KCS1621). g) Poikiloblastic ep1 650 parallel to Ss, with a weak internal foliation defined by qtz (KCS1621).

655

660


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Figure 3. Outcrop photos of epidote boudins sampled for oxygen isotope thermometry. a) SY1613 (Lotos), b) SY1617 (Delfini), c) SY1618 (Delfini), d) SY1623 (Delfini).


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Figure 4. Photomicrographs of measured quartz inclusions in garnet from Delfini (a) and Raman spectrums of unstrained Herkimer 665 quartz and strained quartz inclusions (b). b) Shown for comparison are Herkimer quartz (red) and quartz inclusion (blue) measurements from Virginia Tech and ETH Zürich. Quartz bands and Ar plasma lines (only VT analyses) are numerically labelled.


26

Figure 5. Comparison of Pinc determined from different quartz bands using hydrostatic calibrations, and by using phonon-mode Grüneisen tensors (strains). Red, blue, and yellow symbols indicate qtz-in-grt (Group), qtz-in-ep1 (Group 2), and qtz-in-ep2 (Group 670 3) results, respectively. Diamonds, squares, and circles indicate 𝑷𝑷𝒊𝒊𝒊𝒊𝒊𝒊

𝟒𝟒𝟒𝟒𝟒𝟒𝒗𝒗𝒗𝒗 𝑷𝑷𝒊𝒊𝒊𝒊𝒊𝒊 𝟏𝟏𝟏𝟏𝟏𝟏, 𝑷𝑷𝒊𝒊𝒊𝒊𝒊𝒊

𝟒𝟒𝟒𝟒𝟒𝟒𝒗𝒗𝒗𝒗 𝑷𝑷𝒊𝒊𝒊𝒊𝒊𝒊 𝟏𝟏𝟐𝟐𝟒𝟒,and 𝑷𝑷𝒊𝒊𝒊𝒊𝒊𝒊

𝟒𝟒𝟒𝟒𝟒𝟒𝒗𝒗𝒗𝒗 𝑷𝑷𝒊𝒊𝒊𝒊𝒊𝒊 𝒗𝒗𝒔𝒔𝒔𝒔𝒔𝒔𝒊𝒊𝒊𝒊𝒗𝒗results, respectively.

No border, filled, and open symbols indicate analyses from Kalamisia, Delfini, and Lotos samples, respectively.


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Figure 6. (a) Reference P-T conditions and (b) P-T conditions deduced from elastic thermobarometry and oxygen isotope thermometry superimposed on modeled Cascadia slap-top geotherm (Syracuse et al., 2010). a) Recalculated Ptrap values from Behr 675 et al. (2018) (Syros) and Ashley et al. (2014) (Sifnos) and are shown in blue and red rectangles, respectively. Metamorphic facies are taken from (Peacock, 1993). b) Ptrap from Groups 1, 2, and 3, that reflect peak (qtz-in-garnet), retrograde blueschist-greenschist facies (qtz-in-ep1), and late greenschist facies (qtz-in-ep2) conditions. Solid red, blue, and yellow lines and rectangles are the Ptrap

isomekes (calculated from the mean residual inclusion pressure of each group) and our best-estimate entrapment conditions, respectively. Transparent lines are Ptrap errors (1σ around the mean) for analyses from Delfini samples. Grey box bounds the range 680 of temperatures calculated from oxygen isotope thermometry of quartz-calcite boudin neck precipitates. Metamorphic facies fields (Peacock, 1993): zeolite (ZE), prehnite-pumpellyite (PrP), prehnite-actinolite (PrAc), pumpellyite-actinolite (PA), lawsonite-chlorite (LC), greenschist (GS), lawsonite-blueschist (LB), epidote-blueschist (EB), epidote-amphibolite (EA), amphibolite (AM), eclogite (EC). RSCM = Raman Spectroscopy of Carbonaceous Material (data from Laurent et al., 2018).

685


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Insights from elastic thermobarometry into exhumation of …...Ep2 crystals range from ~ 50 - 300...

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